Download Genesis and evolution of the kaolin

Genesis and evolution of the kaolin-group minerals during the diagenesis and the
beginning of metamorphism
María Dolores Ruiz Cruz
Departamento de Química Inorgánica, Cristalografía y Mineralogía. Facultad de Ciencias.
Universidad de Málaga. 29071 Málaga (Spain).
Introduction
The kaolin group minerals include four members: kaolinite, dickite, nacrite, and the hydrated
analogous, halloysite. Kaolinite, dickite and nacrite show very uniform chemical composition:
46.54 wt.% SiO2, 39.50 wt.% Al2O3 and 13.96 wt.% H2O, corresponding to a formula
Al2Si2O5(OH)4. The formula of halloysite is Al2Si2O5(OH)4.2H2O.
Typical wet chemical analyses of kaolin minerals frequently include small amounts of Fe, Ti,
K, and Mg (Deer et al., 1976), generally due to the presence of impurities such as Fe- and Tioxides and mica. Improvement in instrumentation, especially in electron microscopy associated
to energy dispersive analyses (EDX) has shown that, in most cases, K present in kaolinite
analyses is due to interlayering with other clay minerals (e.g. Lee et al., 1975), whereas some
other cations, such as Fe, Ti or Cr frequently substitutes in minor amounts in the structure (e.g.
Brookins, 1973; Herbillon et al. 1976; Maksinovic et al., 1981; Mestdag et al. 1980). Maxima
amounts of ionic substitutions (Fe =0.44 apfu; Cr=0.56 apfu - atoms per formula unit for
O5(OH)4) have been described in halloysites (Newman and Brown, 1987). It appears that the type
of foreign cations present in the kaolinite structure mainly depends on the bulk-rock composition,
whereas the amount of substitution appears to be controlled by the structure, the disordered
varieties accepting higher amounts of substitutions (Brindley et al., 1986).
The physical and chemical conditions under which the kaolin minerals form are relatively low
pressures and temperatures. These minerals are typical of three main environments: 1)
weathering profiles; 2) hydrothermal alterations; and 3) sedimentary rocks. The most common
parent minerals from which kaolin minerals develop are feldspars and muscovite. The
transformation of potassium feldspar into kaolin minerals occurs according to the equation:
2 KAlSi3O8 + 3 H2O Æ Al2Si2O5(OH)4 + 4 SiO2 + 2 K(OH)
Solubilities of the several chemical species are pH dependent (Mason, 1952). The pH values
of the natural waters normally lie between 4 and 9; alumina is not soluble in this range; silica
solubility increases parallely to the pH and the alkalis and alkaline earth elements are soluble and
mobile. Thus, kaolinite is easily formed and is widespread in soils developed under hot-wet,
intertropical climates (Chamley, 1989). As a consequence, detrital kaolin minerals are important
components of sedimentary rocks deposited near these areas. In addition, kaolin minerals
frequently grow, from the same phases (feldspars and white mica), during the early diagenesis.
The evolution of these minerals during the burial- or tectonic diagenesis is the aim of this
contribution. The accurate study of this evolution requires, however, the knowledge of the main
methods of identification and differentiation among the several kaolin minerals.
42
MARÍA DOLORES RUIZ CRUZ
Methods of study of kaolin minerals
X-ray diffraction and infrared spectroscopy
Structurally, kaolin minerals consist of a sheet of corner-sharing tetrahedra, sharing a plane of
oxygens and hydroxyls (inner hydroxyls) with a sheet of edge-sharing octahedral with every third
site vacant (dioctahedral). This T-O layer has no charge, and interlayer cations are unnecessary to
form the crystal. The T-O layer has two different surfaces: A surface of oxygens and a surface of
hydroxyls (external hydroxyls). This fact determines the nature of the interlayer bonding between
the successive layers, hydrogen bonding, and the need of a relative shift between adjacent layers
= a/3 (Figure 1). The thickness of this elemental layer+interlayer unit is ~7.1 Å (basal spacing),
and permits a rapid and easy identification of this mineral group by X-ray diffraction, especially
when other 7-Å phyllosilicates (serpentine and chlorite) are lacking.
Distinction among the several members of the group requires, however, the detailed analysis
of the non-basal reflections, which reflect the structural differences among the different members
of the group (Table 1).Kaolinite and dickite structures are based on a similar sequence of layers,
and would be identical if trioctahedral. In kaolinite, the vacant octahedral site is in the same place
in all layers (B or C) (Figure 1). In dickite, the vacant octahedral site alternates between B and C
in successive layers, leading to a 2-layer structure. Nacrite structure is based on a different
stacking sequence; initially interpreted as a 6-layer structure, later was defined as a 2-layer
structure in which the conventional X and Y axes are interchanged (Brindley and Brown, 1980).
The true nature of the halloysite structure is not well known (Giese, 1991). These structural
differences are clearly reflected in the X-ray patterns, when one of the members is dominant.
Identification is based on the presence of some diagnostic reflections; the more useful, grouped
by zones, have been summarized in Table 2, and marked in Figure 2.
In addition, the different patterns of kaolinite vary considerably, showing in some cases
sharp, narrow peaks and in other cases, bad-defined broad peaks, according to the degree of
ordering. It is generally observed that the hkl reflections with k=3n are less influenced than those
with k≠3n. In extreme cases, these later peaks lose their identity and originate two-dimensional
modulated bands, similarly to halloysite (Figure 2B).
Other structural differences among the several members of this group affect to the orientation
of the inner and external hydroxyl groups. These differences can be only detected by
spectroscopic methods. The more widely used has been the infrared spectroscopy (FTIR)
although Raman spectroscopy supplies important information. The general features of the OH-
FIGURE 1. Left: [010] view of the elemental structure of the kaolin group minerals. Right: Types
of octahedral positions.
Genesis and evolution of the kaolin-group minerals during the diagenesis and the beginning of metamorphism
43
TABLE 1. Crystallographic data of the kaolin minerals.
System
Space group
Cell parameters
Kaolinite
Triclinic
C1
Dickite
Monoclinic
Cc
Nacrite
Monoclinic
Cc
Halloysite
Monoclinic
Cc
a = 5.156 Å
b = 8.945 Å
c = 7.05 Å
a = 5.138 Å
b = 8.918 Å
c = 14.389 Å
a = 8.908 Å
b = 5.146 Å
c = 15.697 Å
a = 5.14 Å
b = 8.90 Å
c = 14.70 Å
References
α = 91.697º
β = 104.862º
γ = 89.823º
β = 96.74º
Joswig and Drits, 1986
β = 113.7º
Blount et al., 1969
β = 96º
Chukhrov & Zvyagin, 1966
Bish and Von Dreeh, 1988
stretching absorption bands are well established for kaolinite and dickite although some
uncertainties remain concerning nacrite (Farmer, 1976; Russel and Fraser, 1994).A typical
spectrum of kaolinite show four bands, at 3697, 3669, 3652 and 3620 cm-1, whereas the spectra
of dickite and nacrite present only three bands, at slightly different frequencies (Table 3 and
Figure 3). The 3620 cm-1 band has been ascribed to the inner hydroxyls, and the other three (or
two) bands are generally ascribed to vibrations of the external hydroxyls.
FIGURE 2. XRD patterns obtained from unoriented samples. A: Ordered kaolinite from the Campo de
Gibraltar area (Ruiz Cruz and Reyes, 1998, modified). B: Disordered kaolinite from Georgia (standard
KGA-2). C: Ordered dickite from the Campo de Gibraltar area (Ruiz Cruz and Reyes, 1998, modified). D:
Ordered nacrite from the Maláguide Complex (Ruiz Cruz, 1996, modified). In this later pattern the
reflections of dolomite have been deleted.
44
MARÍA DOLORES RUIZ CRUZ
TABLE 2. Main diagnostic reflections (Å) and intensities (I) for kaolinite, dickite and nacrite.
Kaolinite
4.180
4.130
I
50
30
3.842
45
2.558
2.526
2.491
60
40
80
2.338
2.288
2.247
90
80
20
Dickite
4.260
4.120
3.954
3.790
I
8
65
10
55
3.262
2.558
10
35
2.505
50
2.412
2.386
2.324
20
15
95
Nacrite
I
4.130
70
3.476
3.413
20
20
2.432
2.404
60
40
2.321
15
Electron microscopy study
Morphologies, as observed by scanning electron microscopy (SEM), can also be useful in
identification of kaolin minerals. Halloysite shows generally tubular morphologies (Figure 4A)
although spherical particles are also common. Formation of curved layers has been related to the
higher tetrahedral Al for Si substitution, which would originate a mismatch between the
octahedral and the tetrahedral sheets (Giese, 1991).
Kaolinite also shows a variety of morphologies, including platy, pseudohexagonal particles,
booklets and vermicular stacks (Figure 4B). Dickite appears generally as larger blocky particles
or stacks (Figure 4C). Available data for nacrite indicate that this mineral mainly form thin
hexagonal particles (Figure 4D).
Data provided by synthetic kaolin minerals indicate, as previously suspected from natural
occurrences, that morphologies are clearly related to the conditions (mainly fluid/rock ratio, pH
and temperature) and mechanisms of formation. Experimental data indicate that a rapid
precipitation leads to spherical morphologies, whereas slower recrystallization processes
originate platy particles and stacks (Bentabol et al., 2006).
TABLE 3. OH-stretching bands for kaolin-group minerals
Kaolinite
Dickite
Nacrite
Halloysite
υ1
3697
3704
3703
3696
υ2
3669
υ3
3652
3654
3647
3620
υ4
3620
3622
3629
FIGURE 3. Typical FTIR spectra of kaolinite, dickite and
nacrite. Provenance of samples as in Figure 1.
Genesis and evolution of the kaolin-group minerals during the diagenesis and the beginning of metamorphism
45
FIGURE 4. Scanning electron microscopic images showing the most common morphologies of kaolinminerals. A: Tubular halloysite in a sandstone from the Campo de Gibraltar area. B: Vermicular kaolinite
from the Campo de Gibraltar area (Ruiz Cruz and Reyes, 1998, modified). C: Blocky dickite and stacks
from Permo-Triassic sandstones of the Maláguide Complex. D: Thin platy particles of nacrite from
Paleozoic greywackes from the Maláguide Complex (Ruiz Cruz, 1996, modified).
The study by transmission electron microscopy (TEM), the most useful technique for a
precise identification at the nanometer scale, is hindered, in the case of the kaolin minerals, by
the rapid damage of the crystals against the electron beam, which is especially evident in the case
of kaolinite (Figure 5). The selected area electron diffraction patterns (SAED) permit, however, a
rapid differentiation between 1-layer (kaolinite) and 2-layer (dickite or nacrite) structure (Figure
5, A and B, inset).
Evolution of kaolin minerals during the diagenesis
In contrast to the smectite-to-illite dioctahedral sequence, which has been extensively studied
(Merriman and Peacor, 1999 and references therein), the other dioctahedral series, which evolved
FIGURE 5. Lattice-fringe images and SAED patterns of kaolinite from the Campo de Gibraltar area (A) and
dickite from the Sierra Arana Triassic formations (B).
46
MARÍA DOLORES RUIZ CRUZ
from kaolinite toward pyrophyllite, has been less studied. This is because, although kaolin
minerals are common in pelitic sequences, they appear generally in very low amounts. Kaolin
minerals are, however, abundant in two characteristic types of rocks: Fine-grained rocks with
anomalous compositions (tonsteins) and coarse-grained rocks (sandstones and conglomerates).
The term tonstein strictly designates kaolinite-rich claystone beds interbedded with coalbearing strata. These sediments generally extend over large areas and are considered to have a
major volcanic origin. Although kaolinite is the dominant mineral, it generally coexists with
minor amounts of smectite and/or illite. Kaolinite from tonsteins is interpreted as formed during
the early diagenetic alteration of tephra layers under acidic conditions.
We will focus this contribution on the diagenetic processes occurring in sandstones and
associated shales, which have been summarized in Figure 6.
FIGURE 6. Summary of the diagenetic-to-metamorphic parameters (modified from Merriman and Peacor,
1999) characterizing the transformations of kaolin-group minerals during the diagenesis and beginning of
metamorphism. 1: Typical Triassic transition in the North See and in the Iberian and Betic ranges. 2.
Transition observed in strongly deformed shales from El Campo de Gibraltar area (Ruiz Cruz and Reyes,
1998). 3. Transition observed by Shutov et al. (1970) and Ruiz Cruz (1996). 4. Other nacrite reports
(Bühmann, 1988; Buatier et al., 1996). 5. After numerous sources (e.g. Eherenberg and Nadeau, 1989; Ruiz
Cruz and Andreo, 1996a, b; Ruiz Cruz and Sanz de Galdeano, 2005; Ruiz Cruz et al., 2005). 6. Alter
numerous sources (e.g., Merriman and Peacor, 1999).
Genesis and evolution of the kaolin-group minerals during the diagenesis and the beginning of metamorphism
47
Early diagenetic reactions
Sandstones are suitable rocks for the formation of secondary diagenetic minerals because
their high porosity and permeability, which favour the migration of fluids. Most sandstones
contain, in addition to quartz, some detrital feldspars and mica grains, which are favourable
parent phases for kaolinite formation. It is difficult to distinguish unambiguously between detrital
and authigenic clay minerals; nevertheless, detailed optical studies help to this distinction.
Authigenic kaolin minerals generally are pore-lining, pore-filling, fracture-filling, and as
pseudomorphs after previous phases, mainly white mica and feldspars, whereas detrital grains
mainly occur as a dispersed matrix, flocculos and shale clasts.
Although halloysite can form in some sandstones, the most common kaolin mineral is
kaolinite. In most cases its origin is clearly related to leaching of preexisting minerals, which is
favoured at acidic conditions and in the presence of organic mater. During the pH rise subsequent
to dissolution, kaolinite precipitates, since its solubility decreases rapidly as neutral pH values are
reached. Formation of kaolinite through this process is enhanced in continental sandstones, such
as those characteristic of the Permo-Triassic sequences, where kaolinite formed during an early
diagenesis (Ruiz Cruz and Andreo, 1996a). Figure 7A shows the partial kaolintization of a
detrital muscovite grain. In addition, in more altered sandstones, smaller kaolinite crystals fill the
spaces between the grains (Figure 7B). Nevertheless, feldspar and mica dissolution also occur
during the early diagenesis in marine sandstones, such as those deposited in turbiditic sequences
from the Campo de Gibraltar area (Ruiz Cruz, 1994).
Late diagenetic reactions
The evolution of the kaolin minerals at increasing burial- or tectonic depth is different in
pelitic and coarse-grained sequences. In pelitic sequences kaolinite frequently persists until
thermal maturities of 1.9-2.1 %, equivalent to the beginning of the low anchizone (Kisch, 1983)
(Figure 6). In contrast, in the matrix of sandstones, the transformation of kaolinite into dickite has
been reported at approximately 120 ºC, i.e., in the late diagenetic zone, in Triassic sandstones
from the North Sea (Ehrenberg et al., 1993). Similar temperature intervals have been deduced in
Permo-Triassic sandstones from the Betic Cordillera (Ruiz Cruz and Andreo, 1996a) and in
Tertiary sandstones from the Campo de Gibraltar area (Ruiz Cruz, 1994). This different
behaviour can be probably related with the notably higher permeability of sandstone-rich
sequences, which favour the fluid circulation and the process of dissolution-precipitation
responsible of this change.
FIGURE 7. Optical images showing two stages of kaolin-mineral formation in sandstones from the
Maláguide Complex. A: Incipient muscovite-to-kaolinite transformation. B: Kaolinite and dickite porefilling.
48
MARÍA DOLORES RUIZ CRUZ
Although this is an unquestionable pattern, other data clearly indicate that the dickite
formation is not only temperature-dependent, as proposed by Eherenberg et al. (1993). Thus, in
Cretaceous shales from El Campo de Gibraltar area, where kaolinite and dickite fill small
fractures in strongly deformed shales, the isotopic data indicate that the kaoliniteÆdickite
transformation has occurred at lower temperatures (<100ºC), near the transition between the early
and the late diagenetic zones (Ruiz Cruz and Reyes, 1998). Indeed, in this area, dickite coexists
with R0 and R1 illite/smectite mixed-layers. Strain-related dickite has also been described by
Buatier et al. (1997) in Pyrenean thrust-fault zones, also suggesting that the kaolinite-to-dickite
transition is favoured by strain.
Differences observed among the XRD patterns of several size-fractions of sandstones (2-20
µm and <2 µm) and between sandstones and interbedded shales, clearly indicate that the
kaolinite-to-dickite transformation mechanism is different in coarse- and fine-grained rocks (Ruiz
Cruz and Moreno Real, 1993). These authors interpreted these differences as due to the presence
of randomly ordered kaolinite/dickite mixed-layers, which would be intermediate steps in the
kaolinite to dickite transformation.
Whereas the kaolinite-to-dickite transition is widespread in sandstones, the transition
dickiteÆnacrite has been only rarely found. Shutov et al. (1970) recorded nacrite replacing
dickite in deformed veins from Paleozoic and Riphean rocks. Ruiz Cruz (1996) described nacrite
in Paleozoic greywackes underlying dickite-bearing Permo-Triassic rocks. Nacrite grew in small
veins associated to dolomite (Figure 8). Although data from diagenetic terrains are scarce, the
transition kaolinite Æ dickite Æ nacrite, at increasing temperatures has been well documented in
hydrothermal deposits, using isotopic data (Katsumi, 1989).
Nevertheless, thermodynamic determinations clearly indicate that the relative stability of the
several members of the kaolin group does not change significantly with pressure and temperature
over their range of occurrence, kaolinite being the stable phase (de Ligny and Navrotsky, 1999).
Thus, dickite and nacrite are metastable phases, and the described patterns must be interpreted in
terms of kinetics rather than as resulting from changes in the thermodynamically stable
assemblage.
Recent reports of nacrite as well as experimental results agree with this later interpretation.
Thus, at the Lodève basin, nacrite occurs in dolomite cavities associated to barite deposits of
hydrothermal origin (Buatier et al., 1996). Study of the fluid inclusions indicated the presence of
FIGURE 8. Optical microscopic image showing
nacrite associated with dolomite in a Paleozoic
greywacke from the Malaguide Complex (Ruiz
Cruz, 1996, modified).
FIGURE 9. Optical microscopic image showing the
beginning of the illitization process in a PermoTriassic sandstone (Ruiz Cruz and Andreo, 1996a,
modified).
Genesis and evolution of the kaolin-group minerals during the diagenesis and the beginning of metamorphism
49
high-salinity brines, and heating runs indicated temperatures of formation between 80 and 100
ºC. In addition, Bühmann (1988) described authigenic nacrite, which forms a thin layer coating a
carbonaceous shale, originated at ambient temperature. Precipitation from saturated pore
solutions was suggested as the factor determining the precipitation. Our experimental results
(unpublished data) indicate that nacrite formation is favoured by the presence of saturated
solutions, and a rapid precipitation process. At these conditions, nacrite originates metastable
spherical particles similar to those observed in synthetic kaolinite and halloysite, which rapidly
evolves toward platy particles and kaolinite stacks through a slower dissolution-precipitation
process.
All these data indicate that nacrite is not a valid kaolin-mineral indicator of high temperature,
as previously assumed (Hanson et al., 1981). On the contrary, nacrite genesis is temperature–
independent and is more clearly related to the presence of saturated solutions.
Very low-grade metamorphic reactions
The transition kaolin mineral Æ pyrophyllite approximately marks the transition diagenesismetamorphism (Figure 6). This reaction is known many years ago, and has been extensively
studied in natural and experimental systems (Frey, 1987; Bucher and Frey, 1994). Nevertheless,
this reaction only occur in very Al-rich rocks, since, in the presence of K , Mg or Fe, the kaolin
minerals evolve towards other minerals before the temperature necessary for the pyrophyllite
formation was reached. Illitization of kaolin minerals has been frequently described and the
estimated temperatures are in the order of 140 ºC (Ehrenberg and Nadeau, 1989). Thus, in
Permo-Triassic rocks from the Betic and the Iberian ranges, a complete evolution from dickite to
illite can be observed, which is not strictly related to the depth but to the bulk-rock composition.
Although illitization is the dominant process (Figure 9), formation of chlorite and of
illite+chlorite stacks from dickite is also frequently observed (Ruiz Cruz and Andreo, 1996a).
Bulk-rock composition also appears to control other less-known transformations of kaolin
minerals. Thus, in Triassic rocks of the Betic Cordillera, dickite evolves toward sudoite (Ruiz
Cruz and Sanz de Galdeano, 2005), whereas in Carboniferous rocks, dickite and nacrite appear
locally transformed into tosudite (Ruiz Cruz and Andreo, 1996b) (Figure 10). Whereas sudoite is
widespread in rocks with variable grain-size and lithologies, tosudite formation is limited to Alrich rocks, and Al-rich microdomains.
FIGURE 10. Left: TEM image of sudoite formed from dickite in Triassic sandstones from the Sierra Arana
area (Ruiz Cruz and Sanz de Galdeano, 2005). Right: SEM image showing the transformation of nacrite
into Tosudite (Ruiz Cruz and Andreo, 1996b, modified).
50
FIGURE 11. XRD patterns (oriented samples) of
the <2µm size-fractions of two samples from the
Sierra Arana sector. Ms: Muscovite/Illite. Prl:
Pyrophyllite. Dk: Dickite. Qtz: Quartz.
MARÍA DOLORES RUIZ CRUZ
FIGURE 12. Optical microscopic image showing
the dickite to pyrophyllite transformation in a
Triassic sandstone from Sierra Arana (Ruiz Cruz
et al., 2005, modified).
In some Al-richer protolites, the kaolin minerals persist at higher temperatures and originate
pyrophyllite (Figure 12). Generally, temperature controls the dickite/pyrophyllite ratio, and the
Al2O3 content determines the illite/pyrophyllite ratio. In these rocks, dickite, pyrophyllite and
illite coexist in a short temperature interval (Figure 9). The transformation of dickite into mica is
assumed that occurred at high diagenetic conditions. On the contrary, pyrophyllitization of
dickite is assumed to have occurred at incipient metamorphic conditions.
Conclusions and future research
1. The kaolinite Æ dickite Æ nacrite transition is well documented in natural environments
and occurs at increasing temperatures.
2. Thermodynamic and experimental data indicate, however, that the stable phase, at the
common P-T conditions, is kaolinite. Thus, the described transformations are not only
controlled by temperature.
3. Strain and solutions composition appear to be important factors controlling the dickite and
nacrite formation.
4. Futures researches must be focused in natural fluid characterization by means of stable
isotopic and fluid inclusions studies. Parallel synthesis in well-controlled chemical
systems must supply additional information.
Acknowledgements
The author is grateful to Dr. Bentabol, which supplied numerous data about synthetic
kaolinite and to J.M. Garrido for drawing kaolin-mineral structures.
References
Bailey, S.W. (1984). Structures of layer silicates. Pp. 1-123 in: Crystal Structures of Clay
Minerals and their X-ray Identification (G.W. Brindley and G. Brown editors). Mineralogical
Society, London.
Bentabol, M., Ruiz Cruz, M.D., Huertas, F.J., Linares, J. (2006). Clays Clay Miner., 54, 667-677.
Bish, D.L. and Von Dreele, R. (1988). Annual Meeting of the Clay Minerals Society, Michigan.
Blount, A.M., Threadgold, I.M., Bailey, S.W. (1969). Clays Clay Miner., 17, 185-194.
Brindley, G.W., Chih-Chun Kao, Harrison, J.L., Lipsicas, M., Raythatha, R. (1986). Clays Clay
Miner., 34, 239-249.
Genesis and evolution of the kaolin-group minerals during the diagenesis and the beginning of metamorphism
51
Brookins, D.G. (1973). Clays Clay Miner., 21, 421-422.
Brown, G. and Brindley, G.W. (1980). X-ray Diffraction Procedures for Clay Mineral
Identification. Pp. 305-360 in: Crystal Structures of Clay Minerals and their X-ray
Identification. (G.W. Brindley and G. Brown, editors). Mineralogical Society, London.
Buatier, M.D., Deneele, D., Dubois, M., Ptdevin, J-L., Lopez, M. (1996). Eur. J. Mineral., 8, 847852.
Buatier, M.D., Travé, A., Labaume, P., Potdevin, J.L. (1997). Eur. J. Mineral., 9, 875-888.
Bühmann, D. (1988) Clays Clay Miner., 36, 137-140.
Bucher, K. and Frey, M. (1987). Petrogenesis of metamorphic rocks. Springer-Verlag, Berlin,
318 p.
Chamley, H. (1989). Clay sedimentology. Springer-Verlag, Berlin, 623 p.
Chukhorv, F.V. and Zvyagin, B.B. (1966). Proc. Int. Clay Conf. Jerusalen, 11-25.
Deer, W.A., Howie, R.A., Zussman, M.A. (1976) Sheet Silicates. Rock Forming Minerals, vol. 3,
270 p.
Eherenberg, S.N. and Nadeau, P.H. (1989). Clay Miner., 24, 233-253.
Ehrenberg, S.N., Aagaard, P., Wilson, M.J., Fraser, A.R., Duthie, D.M.L. (1993). Clay Miner.,
28, 325-352.
Farmer, V.C. (1974). The layer silicates. Pp. 331-365 in: The Infrared Spectra of Minerals (V.C.
Farmer editor) Mineralogical Society, London.
Frey, M. (1987). Very low-grade metamorphism of clastic sedimentary rocks. Pp. 9-58 in: Lowtemperature metamorphism (M. Frey ed.) Blackie, Glasgow, U.K.
Giese, R.F. (1991). Kaolin minerals: structures and stabilities. Rev. Mineral., 19, 29-66.
Hanson, R.F., Zamora, R., Keller, W.D. (1981). Clays Clay Miner., 29, 451-453.
Herbillon, A.J., Mestdagh, M.M., Vielvoye, L., Derouane, E. (1976). Clay Miner., 11, 201-220.
Joswig, W. and Drits, V.A. (1986). N. Jb. Mineral. Mh., 19-22.
Katsumi, M. (1989). Geochim. Cosmochim. Acta, 53, 2915-2924.
Kisch, H.J. (1983). Mineralogy and petrology of burial diagenesis (burial metamorphism) and
incipient metamorphism in clastic rocks. Pp. 289-493 in: Diagenesis in Sediments and
Sedimentary Rocks (G. Larsen and G.V. Chilingar eds). Elsevier, New York.
Lee, S.Y., Jackson, M.L., Brown, J.L. (1975). Clays Clay Miner., 23, 125-129.
Ligny, D. and Navrotski, A. (1999). Am. Mineral., 84, 506-516.
Maksimovic, Z., White, J.L., Logar, M. (1981). Clays Clay Miner., 29, 213-218.
Mason, B. (1952). Principles of Geochemistry. John Wiley and Sons, New York, 274 p.
Merriman R.J. and Peacor D.R. (1999). Very low-grade metapelites: mineralogy, microfabrics
and measuring reaction progress. Pp. 10-60 In: Low-grade metamorphism (M. Frey and D.
Robinson eds.) Blackwell Science, Oxford, U.K.
Mestdagh, M.M., Vielvoye, L., Herbillon, A.J. (1980) Clay Miner., 15, 1-13.
Newman, A.C.D. and Brown, G. (1987) The Chemical Constitution of Clays. Pp. 1-128 in:
Chemistry of Clays and Clay Minerals. Mineralogical Society, London.
Ruiz Cruz, M.D. (1994). Clay Miner., 29, 93-104.
Ruiz Cruz, M.D. (1996). Clays Clay Miner., 44, 357-369.
Ruiz Cruz, M.D. and Moreno Real, L. (1993). Clays Clay Miner., 41, 570-579.
Ruiz Cruz, M.D. and Andreo, B. (1996a). Clay Miner., 31, 133-152.
Ruiz Cruz, M.D. and Andreo, B. (1996b). Eur. J. Mineral., 8, 1391-1399.
Ruiz Cruz, M.D. and Reyes, E. (1998). Geochem., 13, 95-104.
Ruiz Cruz, M.D., Sanz de Galdeano, C., Lázaro, C. (2005). Eur. J. Mineral., 17, 81-91.
Ruiz Cruz, M.D. and Sanz de Galdeano, C. (2005). Clays Clay Miner., 53, 639-652.
Russell, J.D. and Fraser, A.R. (1994). Infrared methods. Pp. 11-67 in: Clay Mineralogy:
Spectroscopic and chemical determinative methods (M.J. Wilson editor) Chapman and Hall,
London.
52
MARÍA DOLORES RUIZ CRUZ
Shing, B. and Gilkes, R.J. (1991). Clays Clay Miner., 39, 571-579.
Shutov, V.D., Aleksandrove, A.V., Losievskaya, S.A. (1970). Sedimentology, 15, 69-82.